Carmack (1977), the summertime mean temperature of shelf water in the Pacific Sector of the Southern Ocean (-1,03°C) is somewhat warmer than that in the other two sectors; the mean salinity is 34.46%o. Fig. 3.10, adapted from Ainley and Jacobs (1981), shows vertical sections of temperature and salinity across the Ross Sea Shelf near 170°W, and illustrates several of the water masses described below.

The densest water mass in the Southern Ocean is High Salinity Shelf Water observed over the broad continental shelf in the southwestern Ross Sea. Its salinity increases with depth from 34.75'7c.c to 35.00%o or greater (Jacobs et al., 1970). Since this water mass is nearly isothermal at the surface freezing point (approximately -1.9°C), it is probably formed during winter by freezing and evaporation in the leads and polynyas in the western Ross Sea along the Victoria Land Coast (Jacobs et al., 1985). High Salinity Shelf Water appears with salinities greater than 34.80%o in Fig. 3.10. It is presumed that most of this dense water is dynamically constrained to circulate in a clockwise gyre over the shelf (Jacobs et al., 1970). That portion which escapes from the shelf contributes to the formation of the high salinity variety of bottom water.

Shelf waters with somewhat lower salinities (34.60%o-34.75%o) are still dense enough to contribute to the formation of bottom water. Such shelf waters are

Fig. 3.10. Distribution of temperature (top panel) and salinity (bottom panel) constructed from stations occupied along approximately 170°W across the Ross Sea continental shelf in austral summer. The barrier of the Ross Ice Shelf is indicated at upper right in each panel. Adapted from Ainley and Jacobs (1981).

Fig. 3.10. Distribution of temperature (top panel) and salinity (bottom panel) constructed from stations occupied along approximately 170°W across the Ross Sea continental shelf in austral summer. The barrier of the Ross Ice Shelf is indicated at upper right in each panel. Adapted from Ainley and Jacobs (1981).

present within the eastern Ross Sea (Jacobs et al., 1979) and along the Adelie Coast near 140°E (Gordon and Tchernia, 1972).

Most of the continental shelf in the Pacific Sector is occupied by water with salinity lower than 34.60%o (Carmack, 1977) which is referred to as Low Salinity Shelf Water. While mixtures containing this water are not dense enough to sink to abyssal depths, they may sink along the continental slope and spread at intermediate depths (Carmack and Killworth, 1978; Killworth, 1983) thereby ventilating the water column.

About half of the continental shelf in the Ross Sea is covered by the thick permanent floating glacial ice of the Ross Ice Shelf. Some of the subsurface water observed over the continental shelf is colder than the sea surface freezing point and therefore could not have formed at the sea surface. Since the freezing point of seawater decreases with increasing pressure, this very cold water must be cooled at depth by interaction with the ice shelf. This water is therefore called Ice Shelf Water.

Two distinct layers of Ice Shelf Water have been observed in the Ross Sea (Jacobs et al., 1979). Shallow Ice Shelf Water with salinity less than 34.60%o is seen emerging from the edge of the ice barrier at depths between 50 m and 250 m. Relatively warm water originating from Circumpolar Deep Water penetrates over the continental shelf and beneath the Ross Ice Shelf at approximately 175°W (Jacobs et al., 1979; Pillsbury and Jacobs, 1985). Fig. 3.10 shows a tongue of modified Circumpolar Deep Water with temperatures greater than -1.0°C penetrating the shelf above 250 m. Shallow Ice Shelf Water presumably forms as a result of interaction of this warm water with the ice shelf (Jacobs et al., 1985).

Deep Ice Shelf Water is observed at depths between 300 m and 600 m offshore of the ice shelf. Its salinity is greater than 34.60%o and in Fig. 3.10 appears as a tongue of water colder than -2.0°C near 500 m. MacAyeal (1984) has described the likely formation mechanism of Deep Ice Shelf Water. From its source region in the northwestern Ross Sea, High Salinity Shelf Water flows to the southwest along the western margin in a deep depression under the ice shelf. In the southwestern Ross Sea, the ice shelf is thick and the water column below is relatively thin. Tidal currents are strongest in such regions and capable of supplying the mechanical energy necessary to allow the High Salinity Shelf Water to penetrate the thin layer of melt water which separates it from the base of the ice shelf. Subsequent basal melting reduces the salinity and results in the Deep Ice Shelf Water observed further north.

The circulation on the Ross Sea Shelf has been inferred from density distributions supplemented by some bottom photographic evidence and a few direct current measurements (e.g., Pillsbury and Jacobs, 1985). It is thought that the circulation on the shelf is clockwise with westerly flow along the edge of the ice shelf and eastward flow along the shelf break (Jacobs et al., 1970; Ainley and Jacobs, 1981). The salinity field (which is generally parallel to the density field) supports this view if it is assumed that currents at the shelf floor are negligible (Fig. 3.10).

Waters of the continental shelf regime are separated from those of the oceanic regime by a wedge of relatively fresh water at the shelf break (Fig. 3.10). Ainley and Jacobs (1981) referred to this feature as the Slope Front. This separates the eastward-flowing northern limb of the Ross Sea Shelf circulation from the westward flow of the Circumpolar Deep Water. Gill (1973) noted a similar V-shaped salinity distribution in the southern Weddell Sea, but interpreted the flow field differently: he assumed that there is a substantial westward bottom flow at the edge of the shelf and concluded that both shelf and slope waters are flowing westward. Direct current measurements are needed to resolve the flow in these two regions.

Antarctic Bottom Water

Although some investigators (Gordon, 1978; Killworth, 1983) have noted the potential importance of open ocean convection in bottom water formation, most Antarctic Bottom Water is believed to form near the Antarctic continental margins. There, very cold, dense waters (0 < -1.8°C, S > 34.60%o) of the continental shelf mix with components of Circumpolar Deep Water available near the shelf break and over the continental slope. Due to its high density, this mixture tends to sink to the ocean floor along the continental slope but, because of the influence of the earth's rotation, it is deflected to the left and flows approximately parallel to the isobaths as a contour current. When the dense water finally reaches the ocean floor, it spreads laterally along the bottom of the basins adjacent to Antarctica. Exchange between basins is restricted to deep gaps and fracture zones in the submarine ridge system.

Because of its components from the waters over the continental shelf, newly formed Antarctic Bottom Water is characterized by a maximum in dissolved oxygen and minima in potential temperature, silicate, phosphate, and nitrate. The salinity of bottom water varies but, with the exception of that formed in the northwestern Ross Sea, is characterized by a bottom minimum in salinity. Circumpolar Deep Water, which overlies the bottom water, is also identified by minima in phosphate and nitrate, so there must be an induced nutrient maximum between newly formed Antarctic Bottom Water and Circumpolar Deep Water. We will refer to the layer of maximum phosphate, nitrate, and silicate as "old" bottom water. Away from source regions, the newly-formed bottom water is absent and the layer of old bottom water intersects -the bottom. Even further away, the old bottom water is absent and Circumpolar Deep Water intersects the bottom.

The contrast between new and old bottom water is strongest near source regions and is particularly evident in vertical profiles of silicate presented by Edmond et al. (1979) from various sites in the Southern Ocean and in the vertical sections of phosphate and silicate presented by Reid et al. (1977, fig. 6) for the Atlantic Sector. Similar stratification is observed south of New Zealand near the Ross Sea source of bottom water in vertical sections of silicate and nitrate (Craig et al., 1981, plates 83 and 85). Fig. 3.11 shows profiles of properties in the Ross Sea Gyre from a station at 73°S, 159°W. Although the phosphate and nitrate profiles are plotted on an expanded scale that reveals the precision of the measurements, the overall trend is clear: the near-bottom minima in potential temperature and

Fig. 3.11. Vertical profiles of salinity, potential temperature, silicate, oxygen, phosphate, and nitrate obtained at approximately 73°S, 159°W on 15 February 1972 by NORTHWIND (Station 18). Bottom minima in silicate, phosphate, and nitrate indicate presence of newly formed bottom water. Much of the small scale structure in the phosphate and nitrate profiles is not real, but an artifact of the resolution with which the data were recorded.

Fig. 3.11. Vertical profiles of salinity, potential temperature, silicate, oxygen, phosphate, and nitrate obtained at approximately 73°S, 159°W on 15 February 1972 by NORTHWIND (Station 18). Bottom minima in silicate, phosphate, and nitrate indicate presence of newly formed bottom water. Much of the small scale structure in the phosphate and nitrate profiles is not real, but an artifact of the resolution with which the data were recorded.

and nutrients and the maxima in oxygen and salinity show newly formed high-salinity bottom water. The overlying nutrient maxima and relative salinity minima are probably a mixture of old bottom waters from various sources. Further east, new bottom water appears as a bottom minimum in silicate at stations 414 and 415 in Fig. 3.6d.

While it is widely believed that the largest source by volume of Antarctic Bottom Water is in the Weddell Sea, several other source regions have been identified on the basis of water properties that are sufficiently distinct from surrounding bottom waters to indicate local formation. The properties exhibited by newly-formed bottom water reflect the properties of the constituent water masses available at the source. There is speculation that the importance of some sources of Antarctic Bottom Water are not fully appreciated because the waters formed at these sources are essentially indistinguishable from those originating in the Weddell Sea (e.g., Mosby, 1968; Carmack and Killworth, 1978; Jacobs et al., 1985). Within the Pacific Sector, two identifiable source regions are the Ross Sea (Gordon, 1966; Jacobs et al., 1970) and the Adelie Coast between 130°E and 150°E (Gordon and Tchernia, 1972).

Throughout most of the Pacific Sector, salinity values decrease with depth from the salinity maximum (> 34.73%o) of the Lower Circumpolar Deep Water to a minimum (34.69-34.72%o) at the bottom. The northwestern Ross Sea is anomalous due to the introduction of a high-salinity (> 34.72%0) bottom water formed along the continental slope in the western Ross Sea from a mixture of High Salinity Shelf Water and Circumpolar Deep Water (Jacobs et al., 1970). Between the salinity maximum associated with Lower Circumpolar Deep Water and that associated with high-salinity bottom water is an intermediate salinity minimum (< 34.71 %o) which probably represents the upper limit of influence of the high-salinity bottom water. However, Carmack and Killworth (1978) presented evidence to suggest that this deep salinity minimum may be locally reinforced by plumes of low-salinity water flowing down the continental slope.

The high-salinity bottom water initially flows westward towards the Balleny Islands (67°S, 160°E) where it splits into two branches. One branch continues westward, south of the Balleny Islands, into the Australian-Antarctic Basin south of the Southeast Indian Ridge; the other turns to the north, east of the islands, and heads towards the Pacific-Antarctic Ridge where it presumably turns eastward into the Southeastern Pacific Basin (Carmack, 1977). The small areal extent of this high-salinity bottom water suggests that it has a low formation rate and it is quickly diluted by mixing with surrounding waters (Gordon, 1972b).

Anomalously low temperature and salinity and high oxygen values observed off the Adelie Coast (67°S, 140°E) identify this as a source region for bottom water (Gordon and Tchernia, 1972; Carmack, 1977). This Adelie Coast Bottom Water has significant influence on the bottom property distribution within the Australian-Antarctic Basin (Rodman and Gordon, 1982). The branch of high-salinity bottom water advecting westward from the Ross Sea experiences a decrease in salinity and an increase in oxygen due to mixing with Adelie Coast Bottom Water (Gordon and Tchernia, 1972).

Jacobs et al. (1985) reported that a low-salinity bottom water is formed over the continental slope in the eastern Ross Sea from a mixture of Circumpolar Deep Water and the local shelf water (containing some Ice Shelf Water), which is less saline than that in the western Ross Sea. This bottom water presumably contributes to the rapid dilution of the high-salinity variety. On the other hand, Carmack (1977) speculated that it is low-salinity bottom water, primarily of Weddell Sea origin, flowing from the west into the Southeastern Pacific Basin that dilutes the high-salinity Ross Sea bottom water.

As noted earlier, water parcels tend to flow along surfaces of constant potential density, and these surfaces slope down toward the north. Except for a few deep fracture zones, the Southeast Indian Ridge, Pacific-Antarctic Ridge and East Pacific Rise are continuous at depths greater than about 4,000 m. Density values associated with Antarctic Bottom Water do not occur north of this ridge system, so Antarctic Bottom Water is confined to the Australian-Antarctic and Southeastern Pacific Basins. North of this ridge line, the bottom waters are a mixture of Antarctic Bottom Water and Lower Circumpolar Deep Water (Schlemmer, 1978). The volume of Antarctic Bottom Water in the Pacific is about five times less than in the Atlantic or Indian Ocean Basins (Carmack, 1977).

The mixture of Antarctic Bottom Water and Lower Circumpolar Deep Water that comprises the abyssal waters of the southern Tasman Basin come from the Australian-Antarctic Basin through the Balleny Fracture Zone near 155°E. It appears to flow in a clockwise loop through the Tasman Basin and then eastward through gaps in the Macquarie Ridge and into the Southwestern Pacific Basin (Rodman and Gordon, 1982).

From theoretical considerations, Stommel (1957) predicted equatorward flowing bottom boundary currents along the western sides of ocean basins. The SCORPIO section along 43°S (Warren, 1973) shows the western intensification of the bottom flow against the Chatham Rise, and the influence of Lower Circumpolar Deep Water is clear from the salinity maximum between 3,000 and 4,000 m. Northward abyssal flow extends some 1,000 km east of the intense western boundary jet, and transport within the northward flow is estimated to be about 20 Sv (Warren, 1973, 1976). Similar deep western boundary currents have been observed to exhibit variable, and occasionally quite high, velocities that are capable of significant sediment transport (Londsdale and Smith, 1980; Glasby, 1983; Hollister et al., 1984; Carter and Mitchell, 1987).

The northward-flowing bottom boundary current follows the western margin of the Southwestern Pacific Basin along the Tonga-Kermadec Ridge to about 30°S, where it splits into two parts (Reid, 1986). The northern branch extends to the northern end of the basin near 10°S. Here, the salinity and silicate maxima of Lower Circumpolar Deep Water intersect the bottom, and there is no remaining trace of Antarctic Bottom Water (Mantyla and Reid, 1983). An eastward branch of the abyssal flow extends to the East Pacific Rise before turning to the north (Reid, 1986). The inferred flow pattern in the northern part of the Southwestern Pacific Basin is counter-clockwise, similar to the overlying Subtropical Gyre.

Bottom property distributions suggest a clockwise circulation in both the Australian-Antarctic Basin and the Southeastern Pacific Basin (Gordon, 1972b; Rodman and Gordon, 1982; Mantyla and Reid, 1983; Reid, 1986). The relatively high bottom silicate and low oxygen content of waters in the northern and eastern portions of the Southeastern Pacific Basin suggest long residence times and therefore sluggish flow. Newly-formed Ross Sea bottom waters seem to be confined to the southwestern portion of the basin. Vertical sections across Drake Passage (Sievers and Nowlin, 1984) show that the densest waters flowing from the Pacific to the Atlantic are the older, high silicate bottom waters of the Southeastern Pacific Basin.

Circumpolar Deep Water

By far the most voluminous water mass in the Pacific Sector of the Southern Ocean is Circumpolar Deep Water (Carmack, 1977). It is this water mass which occupies the deep and bottom layers throughout most of the Pacific Ocean. Since Circumpolar Deep Water is continuously flowing around the Antarctic Continent, its properties are relatively homogeneous, but it can be divided into two layers (Gordon, 1967b) that derive their distinguishing properties from two different northern sources.

Lower Circumpolar Deep Water is characterized by a deep maximum in salinity and minima in phosphate and nitrate. These properties are derived from North Atlantic Deep Water that is entrained into the eastward-flowing ACC in the southwestern Atlantic (Reid et al., 1977). Although the extrema that mark the core of Lower Circumpolar Deep Water erode slowly downstream because of both lateral and vertical mixing (Reid and Lynn, 1971), the core is still clearly observed in the flow from the Pacific through Drake Passage (Sievers and Nowlin, 1984). The core of Lower Circumpolar Deep Water rises from depths greater than 3,000 m north of the Subantarctic Front to depths less than 1,000 m south of the Polar Front (Fig. 3.6b). Within the Pacific Sector, salinity values at the core decrease from 34.76%o near the Macquarie Ridge south of New Zealand to 34.73%o in Drake Passage (Fig.3.12). Corresponding phosphate values increase from 2.0 to 2.2 (Ag-atoms/1 and nitrate values increase from 31 to 32 (ig-atoms/1 (Bainbridge, 1981; Craig et al., 1981). The remarkable persistence of these extrema suggests that the distributions of properties deep within the ACC are determined largely by advection rather than diffusion. The highest salinity within the salinity maximum core layer in the Pacific is found close to the position of the Polar Front (cf. Figs 3.7 and 3.12), and has been used by Gordon (1972a) to mark both the axis of spreading of Lower Circumpolar Deep Water and the axis of flow of the ACC. The patchiness in the distribution of salinity within this core layer (e.g., Fig. 3.12) is a commonly reported phenomenon (e.g., Gordon, 1975; Jacobs and Georgi, 1977). Although most of this patchiness is probably an artifact of inadequate sampling, the known mesoscale variability in the ACC suggests that some of the patchiness is real.

Besides the general eastward flow of the ACC, other features of the deep circulation can be inferred from the distribution shown in Fig. 3.12. The northward protrusion of high salinities east of New Zealand is indicative of northward flow of Lower Circumpolar Deep Water in this region. The lower salinity tongue (< 34.74%o) extending northeastward along the southern flank of the Pacific-

Fig. 3.12. Distribution of salinity (7oo) on the surface defined by the deep salinity maximum. Adapted from Gordon and Molinelli (1982).

Antarctic Ridge is also consistent with a topographically steered clockwise gyre north of the Ross Sea. The salinity distribution in Fig. 3.12 suggests that the gyre may extend eastward to Drake Passage. This is consistent wih the shallow and intermediate tracer patterns shown by Reid (1986). At depths greater than 3,000 m (e.g., below this core layer), however, his adjusted flow maps show separate gyres north of the Ross Sea and east of the East Pacific Rise.

Upper Circumpolar Deep Water is characterized by a minimum in dissolved oxygen concentration and relative maxima in phosphate and nitrate. The contrasting extrema between Upper and Lower Circumpolar Deep Water make phosphate and nitrate particularly useful in distinguishing these two deep layers. The properties of Upper Circumpolar Deep Water are entrained into the ACC through lateral exchange with the low-oxygen deep waters in the Indian and Pacific Oceans (Callahan, 1972). The lowest oxygen values within this core layer (< 3 ml/1) are observed in the southeastern Pacific off the coast of Chile (Gordon and Molinelli, 1982). In Drake Passage, just downstream from this source, the oxygen minimum is still less than 4.0 ml/1, the phosphate maximum is greater than 2.4 |j.g-atoms/l and the nitrate maximum is greater than 35 fig-atoms/1 (Sievers and Nowlin, 1984).

The core of Upper Circumpolar Deep Water rises from depths greater than 1,800 m in the Subantarctic Zone to depths less than 400 m in the Antarctic Zone (Fig. 3.6c). In the Subantarctic Zone, this core layer is overlain by Antarctic Intermediate Water. As Upper Circumpolar Deep Water shoals south of the Polar Front, its low oxygen, high phosphate, and high nitrate core begins to erode from above due to vertical mixing with the overlying Antarctic Surface Water, which has relatively high oxygen and low nutrient values. This is reflected in the higher oxygen concentrations and density within the oxygen minimum layer south of the Polar Front as compared to values from the core of the Upper Circumpolar Deep Water north of the Polar Front. Since the Antarctic Surface Water is relatively cold, Upper Circumpolar Deep Water is characterized by an induced temperature maximum within the Antarctic Zone (Fig. 3.6a).

Fig. 3.11 shows that, within the Ross Sea Gyre, the oxygen minimum occurs at approximately the same depth as the salinity maximum, phosphate minimum, and nitrate minimum of Lower Circumpolar Deep Water. Within the Ross Sea Gyre, therefore, virtually all of the Upper Circumpolar Deep Water has shoaled and been eroded away through interaction with Antarctic Surface Water and the water column consists almost entirely of Lower Circumpolar Deep Water and Antarctic Bottom Water. The low oxygen at station TH22 (Fig. 3.6c) occurs at a density too great to have originated in Upper Circumpolar Deep Water north of the Polar Front. Similar conditions are observed in the Weddell Gyre, but the oxygen values at the minimum are lower there than those at the same density north of the Polar Front (Whitworth and Nowlin, 1987). They explain this apparent oxygen depletion on an isopycnal surface as being due to biological processes acting over time.

Antarctic Surface Water

The most distinctive feature of the Antarctic Zone is the relatively cold, low-salinity surface layer referred to as Antarctic Surface Water. As noted earlier, the relatively warm, salty Circumpolar Deep Water rises to shallow depths south of the Polar Front and Antarctic Surface Water is formed by modifying this basic water mass. Heat is lost by interaction with the cold atmosphere and ice, and fresh water is added in the form of precipitation and runoff of ice from the continent. Superimposed on this mean flux of heat and salt through the Antarctic Surface Water layer is a pronounced annual oscillation. Winter cooling and sea ice formation induce vertical convection to produce a winter mixed layer down to the permanent pycnocline. During summer, warming and ice melt create a thin, low density surface layer overlying the residual "winter water" identified by a pronounced subsurface temperature minimum (Fig. 3.6a). The depth of this minimum ranges from approximately 50 m near the Antarctic continent and in the centre of the Ross Sea Gyre to approximately 200 m at the Polar Front.

Distributions of temperature and salinity on the temperature minimum surface are presented by Gordon and Molinelli (1982). Temperature increases from near freezing (approximately -1.9°C) adjacent to the Antarctic continent to approximately 2°C at the Polar Front. The salinity distribution is more complex. There is a patchy band of relatively low salinity (< 34.00%o) to the south of the northern limit of the temperature minimum core layer. Salinity increases from this band toward both the north and south with highest salinities (> 34.30%o) being observed in the Ross Sea Gyre. The increase toward the north is due to mixing with more saline Subantarctic Zone waters. The southward increase in the salinity of the temperature minimum layer may reflect mixing with upwelled Circumpolar Deep Water in these areas. Another contributing factor may be the meridional difference in the effects of sea ice formation and melting. During autumn, surface freezing begins to dominate melting near the continent and sea ice begins to accumulate. As sea ice forms, brine is ejected raising the salinity of the underlying layer of Antarctic Surface Water. Once formed, some of the ice drifts northward away from the continent and toward warmer waters. As the dominance of freezing over melting advances northward, so does the northern edge of the ice pack. The maximum northward extent of the pack remains south of the Polar Front (Deacon, 1982). During spring, melting begins to dominate freezing and the ice edge retreats toward the south. The net effect of this annual freezing cycle is to fractionate fresh water from the southern portions of the Antarctic Surface Water layer and transport it northward.

Antarctic Intermediate Water

Besides Antarctic Bottom Water, the second major water mass spreading northward from sources in the Southern Ocean is Antarctic Intermediate Water. Unlike Antarctic Bottom Water, whose northward flow is concentrated in deep western boundary currents, Antarctic Intermediate Water diffuses northward in all longitudes with a northward component that may vary, but is significantly slower than the eastward component of the ACC. This water mass is apparently formed within the Polar Frontal Zone where the layer of relatively cold and fresh Antarctic Surface Water flows beneath, and is mixed with, warmer, saltier, and less dense Subantarctic waters. Although the subsurface temperature minimum of Antarctic Surface Water is eroded away by vertical mixing within the Polar Frontal Zone, the low salinity waters above the temperature minimum provide Antarctic Intermediate Water with its identifying characteristic, a salinity minimum (Fig. 3.6b). The properties of Antarctic Intermediate Water vary somewhat within and between the major oceans of the southern hemisphere. This is due in part to regional differences in the properties of the Subantarctic contributions to the water mass. The Antarctic Intermediate Water near its formation region in the Pacific is the warmest (> 5°C) and least dense (a ~ 27.1 mg/cm3) variety (Piola and Georgi, 1982).

In the South Pacific, the salinity minimum core layer is near the sea surface in the vicinity of the Polar Frontal Zone (Fig. 3.6b). In some sections, the minimum can be traced southward across the Polar Frontal Zone to the Polar Front (Sievers and Nowlin, 1984). The core layer deepens rapidly toward the north at the Subantarctic Front, reaches depths of approximately 1,000 m within the subtropical anticyclonic (counter-clockwise) gyre, and can be traced northward to at least 10°S (Craig et al., 1981). The low-salinity surface waters in the south-eastern Pacific off the coast of Chile (Fig. 3.8) provide little contrast with the underlying salinity minimum layer. The core of Antarctic Intermediate Water is therefore rather poorly defined in this region.

Spreading of Antarctic Intermediate Water is strongly influenced by the wind-driven surface circulation. Most of the flow in the Pacific Sector is eastward with the ACC. However, there is a westward-intensified, subtropical anticyclonic gyre over the Southwestern Pacific Basin east of the North Island of New Zealand (Reid, 1965; Piola and Georgi, 1982). At approximately 1,000 m, the gyre is centred along 40°S. Core layer properties attenuate along the flow path because of vertical and lateral mixing. Near the source, the core layer is characterized by salinities less than 34.20%o, temperatures less than 4°C, and dissolved oxygen values greater than 6 ml/1. These values attenuate northward along the eastern limb of the subtropical gyre and westward along the northern limb. Within the centre of the gyre and throughout the Tasman Basin, salinities are greater than 34.40%o, temperatures are greater than 5.5°C, and dissolved oxygen is less than 4.5 ml/1 (Gordon and Molinelli, 1982). The properties of Antarctic Intermediate Water in the southwestern Pacific reflect the entrainment of warm, salty, and low oxygen waters from the Coral Sea (Reid, 1965).

Subantarctic Mode Water

One of the distinguishing characteristics of the Subantarctic Zone is a thick, nearly homogeneous layer (referred to as a pycnostad) near the sea surface just to the north of the Subantarctic Front (McCartney, 1977,1982). The large volume of water contained in this pycnostad emerges as an isolated peak or "mode" in regional volumetric censuses (e.g., Cochrane, 1958; Montgomery, 1958; Worth-

ington, 1981; Piola and Georgi, 1982). Consequently, this water has been named Subantarctic Mode Water (McCartney, 1977). The homogeneity of this layer is apparently the result of deep vertical convection driven by winter cooling at the sea surface. The convective overturning also serves to ventilate this water as evidenced by its high dissolved oxygen content (> 6 ml/1). As noted earlier, properties within the Subantarctic Zone vary from west to east across the Pacific and this variation is reflected in the properties of Subantarctic Mode Water observed in the SCORPIO section along 43°S (Reid, 1973; McCartney, 1982; Piola and Georgi, 1982). The potential density (in a units) of this water increases from 26.9 mg cm-3 in the Tasman Basin to 27.1 mg cm'3 in the Southeastern Pacific Basin and is accompanied by a decrease in temperature from 8.5°C to 5.5°C and a decrease in salinity from 34.62 to 34.25%o.

As Subantarctic Mode Water is advected northward by the subtropical anti-cyclonic gyre, it occupies a layer above the salinity minimum of Antarctic Intermediate Water (Piola and Georgi, 1982) and it retains its identifying characteristics of a pycnostad and an oxygen maximum. Although these features attenuate along the flow path, they can be traced to the northern limb of the subtropical anticyclonic gyre along 15°S (Reid, 1973; McCartney, 1977; Craig et al., 1981). The warmer, saltier, and lighter variety of Subantarctic Mode Water observed in the southwestern Pacific is apparently confined to this part of the ocean (McCartney, 1982). The oxygen maximum core layer descends to a maximum depth of approximately 700 m at about 28°S and then rises to approximately 500 m at 15°S where oxygen values at the core have decreased to less than 4 ml/1 (Craig et al., 1981). In the SCORPIO section along 43°S, there is only a weak and poorly-defined oxygen minimum separating the Subantarctic Mode Water from the high oxygen surface layer. However, at 28°S, the separation is distinct (Reid, 1973). This may explain why the subsurface oxygen maximum of Subantarctic Mode Water is not detected in the vertical section shown in Fig. 3.6c, which extends northward only to 45°S.

Shallow Oxygen Minimum Layer and the Peru-Chile Undercurrent

Throughout much of the South Pacific, the weak oxygen minimum layer overlying the maximum in the Subantarctic Mode Water appears as a relative minimum because of the presence of higher dissolved oxygen concentrations in the surface layer (Reid, 1965, 1973). East of about 160°W, this shallow oxygen minimum layer is also characterized by a salinity maximum induced by the overlying low-salinity surface layer of the southeastern Pacific (Fig. 3.8). However, along the eastern boundary of the South Pacific, the extrema in this layer are not merely induced features. Instead, waters with very low dissolved oxygen values (< 2 ml/1) and relatively high salinities (> 34.40%o) are advected into this layer from the tropical Pacific by the southward-flowing Peru-Chile Undercurrent. The water in this current is clearly distinguished from that further west in this same layer by its high phosphate, nitrate and silicate concentration (Reid, 1973). The nutrient concentrations within the core of this undercurrent are high enough to cause the underlying Subantarctic Mode Water to be locally characterized by an induced minimum in nutrients.

The properties marking the core of the Peru-Chile Undercurrent become shallower and attenuate toward the south, but are still clearly evident in the SCORPIO section at 43°S (Reid, 1973), and can be traced southward to approximately 50°S (Silva and Neshyba, 1979/80). At 43°S, the core is at a depth of approximately 250 m with salinity greater than 34.40%o, oxygen less than 3 ml/1, phosphate greater than 2.6 |Jg-atoms/l, nitrate greater than 30 (ig-atoms/1, and silicate greater than 20 |Jg-atoms/l. At 50°S, the core has shoaled to a depth of approximately 200 m with salinity less than 34.20%o, oxygen greater than 4 ml/1, phosphate less than 2.0 |j.g-atoms/l, nitrate less than 30 |jg-atoms/l, and silicate less than 15 |ig-atoms/l. South of this latitude, the undercurrent loses its identifying characteristics.

Subantarctic Surface Water

Although the properties of the surface waters within the Subantarctic Zone are subject to much spatial and temporal variability due to exposure to the atmosphere, these waters are normally characterized by a surface maximum in temperature and dissolved oxygen and a corresponding minimum in nutrients. There is also a general southward decrease in temperature and an increase in oxygen and nutrients. Salinity in the southeastern Pacific has a somewhat more complex distribution because of the extensive tongue of low-salinity surface water (Deacon, 1977b) that extends westward from the South American coast (Fig. 3.8). In the region covered by this tongue, the surface waters are characterized by a salinity minimum; elsewhere the surface waters are characterized by a salinity maximum (Reid, 1973). The apparent source of this feature is the fresh water input along the southern coast of Chile by heavy precipitation and continental runoff (Silva and Neshyba, 1979/80), but the mechanism for distributing this low-salinity water westward remains unresolved. Throughout this region, the large-scale mean flow, in both the atmosphere and ocean, is toward the east. The area occupied by the tongue coincides with the area in which the West Wind Drift splits into northward and southward flowing branches. However, since salinity in this region increases with depth in the upper layer, the low-salinity tongue cannot be maintained by divergence and upwelling. The hypothesis of westward surface flow to maintain this distribution has been advanced (Deacon, 1977b; Neshyba and Fonseca, 1980), but confirmation by direct measurements is lacking.